1.3.1 Historical basis and basic underlying principles
Since the beginning of the Industrial Revolution in the 18th century, considerable amounts of fossil fuel carbon has been burned resulting in increased atmospheric CO2 concentrations from around 280 ppm reaching over 400 ppm in 2014 (Tans and Keeling, 2015). Increased atmospheric CO2 concentration works in a similar way to a greenhouse and traps heat energy in the earth system, observed as an increase in air temperature (i.e. global warming). Due to the high thermal capacity of water and large volume, a major proportion of the heat energy is transferred to the oceans resulting in increased water temperature, termed ‘ocean warming’.
In addition to the warming effect, around 30% of anthropogenic CO2emitted to the atmosphere ends up in the oceans (Sabine et al., 2004), leading to measurable changes in seawater chemistry.
CO2 dissolves in seawater to form a weak acid which dissociates (Eqn. 1.3), releasing H+and causing a shift in the carbonate system in seawater.
CO2+ H2O−−)−−*H2CO3 −−)−−*H++ HCO3− −−)−−*2 H++ CO32− (1.3) This rapid influx of CO2has been clearly observed in major oceans from the Pacific to the eastern and western Atlantic Oceans (Rhein et al., 2013). This change in seawater chemistry is detectable above seasonal variation over a number of decades (Tans and Keeling, 2015), and is at a rate of increase not observed over the geological history (Hönisch et al., 2012). Continued emission of CO2 is expected to increase atmospheric CO2 concentrations to over 1000 µatm with mean projected increases in surface air temperature of 4◦Cand a decrease of around 0.3 in seawater pH by the year 2100 (Collins et al., 2013; Ciais et al., 2013).
1.3.2 Ocean acidification and warming in the Baltic Sea
In addition to the anthropogenic pressures mentioned in Section 1.2.1. Additionally, the Baltic Sea is recognised as a hotspot for both ocean warming and acidification which may change the baseline ecosystem ecology (Elmgren, 2001). Model-based projections suggest that summer surface seawater temperature will likely further increase by between 2 and 4◦Cby the end of this century (HELCOM, 2013) and average pH decrease of around 0.3 – 0.4 (Fig. 1.5, Omstedt et al.
2012) under ‘Business as usual’ (scenario BAU-A2).
The respective decrease in surface water pH from ocean acidification is more difficult to ac-curately predict than for the open ocean because may be partially negated through changes in riverine alkalinity inputs (Schneider et al., 2015) or extent of anoxia in the deeper basins (Haven-hand, 2012). Indeed, change in pH will not be uniformly distributed as regional differences in alkalinity modulate CO2 uptake and the pH decrease (HELCOM, 2013). Nevertheless, it ap-pears as though the rate of change in the Baltic Sea proper, where the terrestrial influence is less than coastal regions, pH is decreasing at a rate faster than in the open ocean (Fig. 1.5). Baltic Sea surface water temperature has increased in all regions since 1990 on the order of 1◦Cper decade (Lehmann et al., 2011), much higher than projections (see above).
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1995 2000 2005 2010
Year pHNBSat25°C
q q
Baltic Proper (BY5 and BY15) Hawaii Ocean Time−series
8.2 8.4
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pH
Surface-water pH qEastern Gotland Basin)
2060 2080 1960 1980 2000 2020 2040
7.6
ECHAM, A2, Business as usual Average annual pH
Year
(a) (b)
Figure 1.5: (a)Measured pH in Baltic Sea and in tropical North Pacific Ocean (HOT times series). Data adapted from Dore et al. (2009) and The International Council for Exploration of the Sea (2014).
(b)Modelled pH data indicating projected future pH decrease, figure adapted from Omstedt et al. (2012).
1.3.3 Biotic response to ocean acidification and warming
These physical and chemical changes in seawater from ocean acidification and warming af-fect the metabolism and pH-sensitive processes and structures in aquatic organisms. Planktonic organisms are too small for internal temperature regulation thus their metabolism is affected directly by the ambient environmental temperature. A concerted research effort has been made over the past two decades which have shown physiological processes in phytoplankton such as C-fixation, biomineralisation, and enzyme-mediated organic matter degradation are temperature, pH and/or CO2-sensitive as comprehensively summarised in two recent publications (Pörtner et al., 2014; Riebesell and Tortell, 2011). RuBISCO is a key enzyme in the fixation of CO2 which is capable of both carboxylation (C-fixation) and photorespiration (remineralisation) ac-tivity, however diffusion of CO2into the cell alone may not provide sufficient carbon for efficient C-fixation (Raven, 1993). Therefore, many phytoplankton invest in carbon concentrating mech-anisms to increase intracellular CO2 and saturate the RuBISCO enzyme. This increase in CO2 at carboxylating site saturates the enzyme RuBISCO and leads to reduced photorespiration at high CO2 (Raven and Beardall, 2003). Ocean acidification will increase the concentrations of dissolved CO2 and hence diffusion into the cell, leading to potential fertilisation of C-fixation and reducing the need for CCM activity. Some organisms may profit more than others due to different affinities of RuBISCO and the energetic relief that increased CO2availability reducing the need for CCM, indicating likely changes in competitiveness between phytoplankton species (Dutkiewicz et al., 2015; Reinfelder, 2011). Nonetheless any stimulating effect of increased CO2 availability will be mediated concurrently by the effect of pH on cellular pH homeostasis (Taylor et al., 2011; Bach et al., 2011).
For diazotrophic organisms studies have shown that H2 production as a by-product of ni-trogenase activity in the photosynthetic bacterium,Rhodobacter sulfidophilus, is pH dependent (Peng et al., 1987), with peak enzyme activity for N2-fixation between pH 7 and 8.2 (Pham and Burgess, 1993). Furthermore, in hetercystous diazotrophic cyanobacteria, such as N. spumi-gena, N2-fixation rates may be negatively affected by pH due to intercellular transfer of carbon
12 CHAPTER 1. INTRODUCTION and nitrogen between the heterocysts and vegetative cells (Czerny et al., 2009). C-fixation also provides energy for N2-fixation. N2-fixation is also expected to be affected by increased CO2 hence may provide a natural new N source to relieve N limitation and stimulate primary produc-tion. Table 1.2 is a current overview of published studies showing how the observed response of N2-fixation activity both between and within diverse diazotrophic genera to increased CO2 has been far from consistent. How diazotrophic organisms respond to ocean acidification may therefore be dependent on differences in O2exclusion strategies and cell morphology, or due to adaptation of CCMs to the carbonate chemistry of their respective ecological niches as suggested by Eichner et al. (2014).
Table 1.2: Summary of reported responses of biomass normalised N2-fixation rates to increased CO2/decreased pH (OA = ocean acidification) in single-strain isolates of a variety of diazotrophic or-ganisms. HL indicates high light and LL indicates low light conditions and the direction of the arrows correspond to the observed direction of the response.
Genus Isolated from Cell
morphology OA
response OW
response Reference Trichodesmium marine, tropical filamentous,
non-heterocystous
⇔ Mulholland and
Bern-hardt (2005)
⇑ Barcelos e Ramos et al.
(2007)
⇑ ⇔ Hutchins et al. (2007)
⇑ Levitan et al. (2007)
⇑ Kranz et al. (2010)
⇑ Garcia et al. (2011)
+Fe⇔
- Fe⇓ Shi et al. (2012)
⇑ Hutchins et al. (2013)
⇑ Fu et al. (2014)
⇑ Hutchins et al. (2015)
Crocosphaera marine, tropical unicellular ⇑ Fu et al. (2008)
⇑ Hutchins et al. (2013)
HL⇑
LL⇔ Garcia et al. (2013b)
Garcia et al. (2013a)
⇑ Fu et al. (2014)
Cyanothece marine unicellular ⇑ Brauer et al. (2013)
⇔ Eichner et al. (2014)
Calothrix marine symbiotic ⇑ Eichner et al. (2014)
Nodularia brackish filamentous,
heterocystous ⇓ Czerny et al. (2009)
(Baltic Sea) ⇑ Wannicke et al. (2012)
⇔ Karlberg and Wulff
(2013)
⇓ Eichner et al. (2014)
Aphanizomenon freshwater filamentous,
heterocystous ⇓ Yamamoto and Nakahara
(2005)
Additionally, temperature strongly modulates N2-fixation and diazotroph abundances, as pre-viously described in Section 1.1.3. However as optimum temperature for N2-fixation was com-monly below 30◦C, interpretation of these studies in terms of ocean warming is complex there-fore results in Table 1.2 should be interpreted with caution. For example, organisms can also increase the geographical distribution as populations may shift polewards to escape physiolog-ically intolerable warm temperatures due to shifting isotherms as suggested forTrichodesmium by Breitbarth et al. (2007) and indicated in model simulations (Dutkiewicz et al., 2015). Even less is known about the interactive effects of CO2 and temperature. It has been shown that the physiological and growth response of phytoplankton to CO2can be modulated by changing tem-perature in single-strain cultures (Sett et al., 2014; Fu et al., 2007). There are comparatively few studies compared to those investigating the individual effects of each driver. Nutrient concentra-tion and other resource availability such as changing light fields through increased stratificaconcentra-tion may also interplay and complicate projections of the biological response of plankton to ocean acidification and warming (Gunderson et al., 2016).